PUBLICATIONS Tectonics REVIEW ARTICLE 10.1002/2017TC004526 Special Section: An appraisal of Global Continental Crust: Structure and Evolution Key Points: • In African rifts that formed within the last 7 Myr, ~20% of the crust is new magmatic material • In mature rift sectors, the surface area of the rift has doubled in ~ 10 Myr, and intrusive volumes are 30% or more • Plate rupture is achieved through catastrophic magma intrusion events in 15–20 km thick transitional crust

Correspondence to: C. J. Ebinger, [email protected]

Citation: Ebinger, C. J., Keir, D., Bastow, I. D., Whaler, K., Hammond, J. O. S., Ayele, A., … Hautot, S. (2017). Crustal structure of active deformation zones in Africa: Implications for global crustal processes. Tectonics, 36. https://doi.org/ 10.1002/2017TC004526 Received 24 FEB 2017 Accepted 4 SEP 2017 Accepted article online 13 NOV 2017

Crustal Structure of Active Deformation Zones in Africa: Implications for Global Crustal Processes C. J. Ebinger1 , D. Keir2,3 , I. D. Bastow4 , K. Whaler5, J. O. S. Hammond6 M. S. Miller8 , C. Tiberi9 , and S. Hautot10

, A. Ayele7,

1

Department of Earth and Environmental Sciences, Tulane University of Louisiana, New Orleans, LA, USA, 2Dipartimento di Scienze della Terra, Università degli Studi di Firenze, Florence, Italy, 3Ocean and Earth Science, University of Southampton, Southampton, UK, 4Department of Earth Science and Engineering, Imperial College London, London, UK, 5School of GeoSciences, University of Edinburgh, Edinburgh, UK, 6Department of Earth and Planetary Sciences, Birkbeck, University of London, London, UK, 7Institute of Geophysics, Space Science, and Astronomy, Addis Ababa University, Addis Ababa, Ethiopia, 8Research School of Earth Sciences, Australian National University, Canberra, ACT, Australia, 9Géosciences Montpellier, UMR5243-CNRS, Université Montpellier, Montpellier, France, 10Imagir sarl, Brest, France

Abstract The Cenozoic East African rift (EAR), Cameroon Volcanic Line (CVL), and Atlas Mountains formed on the slow-moving African continent, which last experienced orogeny during the Pan-African. We synthesize primarily geophysical data to evaluate the role of magmatism in shaping Africa’s crust. In young magmatic rift zones, melt and volatiles migrate from the asthenosphere to gas-rich magma reservoirs at the Moho, altering crustal composition and reducing strength. Within the southernmost Eastern rift, the crust comprises ~20% new magmatic material ponded in the lower crust and intruded as sills and dikes at shallower depths. In the Main Ethiopian Rift, intrusions comprise 30% of the crust below axial zones of dike-dominated extension. In the incipient rupture zones of the Afar rift, magma intrusions fed from crustal magma chambers beneath segment centers create new columns of mafic crust, as along slow-spreading ridges. Our comparisons suggest that transitional crust, including seaward dipping sequences, is created as progressively smaller screens of continental crust are heated and weakened by magma intrusion into 15–20 km thick crust. In the 30 Ma Recent CVL, which lacks a hot spot age progression, extensional forces are small, inhibiting the creation and rise of magma into the crust. In the Atlas orogen, localized magmatism follows the strike of the Atlas Mountains from the Canary Islands hot spot toward the Alboran Sea. CVL and Atlas magmatism has had minimal impact on crustal structure. Our syntheses show that magma and volatiles are migrating from the asthenosphere through the plates, modifying rheology, and contributing significantly to global carbon and water fluxes.

1. Overview The geological record of the African continent spans three quarters of Earth history. Key events occurred in the Archaean when the West African, Congo, Kaapvaal, Zimbabwe, and Tanzanian cratons formed, followed by Palaeoproterozoic accretion of cratons. Unusually, only the northern and northwestern margins of the African plate have been involved in orogeny since the end Pan-African (~500 Ma), when a Himalayan-scale collision formed along eastern Africa and orogenies ringed the Zimbabwe-Kaapvaal-Congo cratons of southern and western Africa (e.g., Muhongo & Lenoir, 1994; Shackleton, 1986; Van Hinsbergen et al., 2011). Each of these processes has left an indelible signature on Africa’s crust and contributed to the formation and destruction of continental topography.

©2017. The Authors. This is an open access article under the terms of the Creative Commons Attribution License, which permits use, distribution and reproduction in any medium, provided the original work is properly cited.

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Excluding the collisional belts of NW Africa (e.g., Miller & Becker, 2014), the predominant tectonic processes since the Pan-African have been magmatism and extension that have increased the surface area and volume of African crust, as outlined in this paper. Southeastern Africa was affected by rifting and flood magmatism at 184–179 Ma (Ferrar-Karroo flood basalts) prior to the separation of Africa and Antarctica (e.g., Duncan et al., 1997), and large sectors of western Africa experienced flood magmatism at 137–127 Ma (Etendeka-Parana flood basalts) as the south Atlantic opened (e.g., Mohriak et al., 2012; Turner et al., 1994). Central, southern, and eastern Africa were affected by the Mesozoic rifting and breakup of Gondwana that left scars such as the Central African rift system and the widespread extensional basins from Somalia to S. Africa (e.g., Burke, 1996; Veevers et al., 1994). Over the past ~45 Ma, Africa again has experienced kimberlitic to flood magmatism and plateau uplift, followed by diachronous rifting from Egypt south to South Africa and southwest

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Figure 1. Topography of Africa with major plates (after Bird, 2003). Landforms and tectonic features labeled. White triangles are temporary seismic networks, whereas black solid shapes are permanent seismic stations. Stations used in controlled source experiments shown in Figures 5 and 7. Boxes enclose regions with crustal thickness information constrained from receiver functions in Figures 2 and 3. Subregions are explored in more detail in subsequent figures. SW = Southwestern rift, WR = Western rift, ER = Eastern rift, MER = Main Ethiopian Rift, Rovuma mp = Rovuma microplate, Victoria mp = Victoria microplate.

through Botswana (e.g., Ebinger & Scholz, 2012) (Figure 1). The Tertiary rift basins and magmatic provinces include all stages and styles in the evolution of cratonic rift zones: incipient rifting in the Okavango to incipient seafloor spreading in the Afar depression, amagmatic sectors in the Western and Southwestern rift to the large igneous province in Ethiopia-Yemen, and sectors that transect Paleozoic lithosphere of variable thickness and composition (e.g., Corti, 2009; Ebinger & Scholz, 2012; Rooney et al., 2017) (Figures 1–4). Oligocene-Recent magmatism with little extension modifies crustal structure beneath West-Central Africa along the Cameroon Volcanic Line (CVL) (Figure 3). Perhaps owing to the very slow movement of the African plate since collision with Europe initiated at ~40 Ma, domal uplift and magmatism without faulting also characterize northern Africa (e.g., Hoggar-Tibesti; Bond, 1978; Brown & Girdler, 1980; Rosenbaum & Lister, 2004) (Figure 1). Over the past few decades, multiple geoscientific experiments have yielded fundamental new constraints on the structure, composition, and evolution of the crust beneath extensional basin systems in Africa and between Africa and Arabia (e.g., Berckhemer et al., 1975; Prodehl & Mechie, 1991a, 1991b; Tokam et al., 2010; Wölbern et al., 2010). Controlled source seismic experiments provide absolute velocity control and details along largely 2-D profiles of crust and, where seismic sources are large, the upper mantle (e.g.,

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Figure 2. Crustal thickness in East Africa constrained using receiver function analysis. Boxes in Figure 2a enclose area of Figure 8 and Figure 11. Box in Figure 2b encloses area shown in Figure 5. Data are sourced from (a) Ethiopia-Yemen plateau regions and the Gulf of Aden Island of Socotra (Ahmed et al., 2014, 2013; Dugda & Nyblade, 2006; Hammond et al., 2011; Stuart et al., 2006), (b) East Africa (Dugda et al., 2005; Hodgson et al., 2017; Plasman et al., 2017; Tugume et al., 2012; Wölbern et al., 2010), and (c) Southern Africa (Nguuri et al., 2001; Kgaswane et al., 2009).

Keller et al., 1994; Maguire et al., 2006; Ruegg, 1975; Stuart et al., 1985). Receiver functions provide constraints on velocity contrasts at the Moho and, in some areas, intracrustal and intramantle reflectors, beneath each permanent or temporary seismic station (e.g., Gao et al., 2013; Hammond et al., 2011; Hodgson et al., 2017; Plasman et al., 2017) (e.g., Figures 2 and 3). Crustal tomography provides a 3-D image of the velocity structure and serves to connect velocity variations between the 2-D controlled source profiles and receiver function point measurements. Ambient noise and arrival time tomography methods, however, rarely image the lower crust owing to sparse ray coverage from local earthquakes and array aperture for ambient noise (e.g., Accardo et al., 2017; Daly et al., 2008; Kim et al., 2012; Korostelev et al., 2015). Magnetotelluric data provide 2-D profiles and point measurements of crust and upper mantle electrical resistivity

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Figure 3. Crustal thickness in (a) the Atlas orogen (Africa-Eurasia collision) and (b) Cameroon Volcanic Line constrained using receiver function analysis. Z-Z0 denotes location of profile shown in Figure 14. Data from Figure 3a are Cooper and Miller (2014), Jessell et al. (2016), Mancilla and Diaz (2012); Miller and Becker (2014), and Spieker et al. (2014) and Figure 3b are sourced from Tokam et al. (2010) and Gallacher and Bastow (2012).

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Figure 4. Topography of Africa with major plates, and 1976–2016 seismicity from the NEIC catalogue, and Holocene to Recent volcanoes from the Global Volcanism Program (http://volcano.si.edu/). ER = Eastern rift, WR = Western rift, SW = Southwestern arm, Rovuma mp = Rovuma microplate, Victoria mp = Victoria microplate. Large Archaean cratons with deep roots labeled; small Tanzania craton lies between the ER and WR. Lower crustal earthquakes with depths greater than 15 km are highlighted in magenta. Boxes enclose regions with crustal thickness information constrained from receiver functions in Figures 2 and 3.

properties, including resistivity versus depth and azimuthal anisotropy of electrical conductivity (e.g., Selway et al., 2014; Whaler & Hautot, 2006). Magnetotelluric data are more sensitive to the presence of melt than seismic methods. Although models of gravity anomalies are highly nonunique, when constrained or jointly inverted with independent 2-D and 3-D data sets, the spatial patterns of gravity anomalies provide additional information on 3-D density variations (e.g., Roecker et al., 2017; Tiberi et al., 2005). The combination of magnetotelluric, seismic, and gravity data in separate and joint inversions enables tighter constraints on material properties and, in some instances, provides new insights into the distinction between magmatic fluids, aqueous fluids, and volatiles within the crust and their role in crustal deformation. This contribution synthesizes our current understanding of crustal modification in tectonically active cratonic rift zones and mountain belts and lays out a road map for future studies of African rift and orogenic zones. Several large-scale crust and mantle imaging experiments have acquired promising data sets in incipient and weakly extended, magma-poor rift sectors (e.g., Gao et al., 2013; Hodgson et al., 2017; Shillington et al., 2016), but a synthesis of crustal structure of early stage rifts is premature. Our focus is to characterize crustal structure in zones of active deformation in the East African rift, Cameroon Volcanic Line (CVL), and

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Atlas Mountains. The review of along-axis variations in the East African rift affords the opportunity to compare and contrast magma-rich rift sectors, providing insights on the evolution of the continental crust in response to stretching and magmatism. This synthesis and comparison offers new insights into the nature of “transitional” crust beneath late-stage rifts (e.g., Fuis & Mooney, 1990; Persaud et al., 2016) and passive margins worldwide (e.g., Skogseid et al., 2000; Van Avendonk et al., 2009). Comparison and contrast with the magmatically modified crust beneath the Atlas Mountains and CVL add to our understanding of fluid and volatile migration through the African crust. Section 2 provides a broad-brush summary of the geodynamical context of active deformation in Africa. Section 3 outlines the role of magma intrusion during the first 5–7 Myr of rifting in cratonic lithosphere, using examples from the southern sector of the Eastern (Gregory) rift. Sections 4 and 5 address the role of magmatism and thinning in the formation of crust transitional between continental and oceanic. The Horn of Africa is one of few areas worldwide where this transition is occurring, and comparisons with the well-studied Salton Trough and northern Gulf of California (e.g., Fuis & Mooney, 1990; Persaud et al., 2016) inform our understanding of magmatic modification along passive margins worldwide. In section 6, we journey across the continent, where warm asthenosphere rises beneath the eastern edge of the Congo craton and the southern margin of a preexisting Mesozoic rift zone. This linear belt of Cenozoic eruptive volcanic centers, neither the CVL, exhibits the age progression, mechanical stretching, magmatic modification displayed in parts of East Africa nor along linear volcanic tracks in oceanic plates. Here too, volatiles have likely played a central role in magmatic modification of Africa’s crust. Section 7 presents a review of crustal structure and its relation to mantle dynamics beneath the Atlas Mountains where compressional tectonics may inhibit upward migration of melt, leading to the localized, linear trend of magmatism linked instead to the Canary Island hot spot. This integration of constraints on crustal structure beneath zones of active rifting and orogenesis in Africa shows that magma and volatiles are migrating from the asthenosphere through the plates, modifying lithospheric rheology and significantly contributing to global carbon and water fluxes.

2. Geodynamic Settings The slow-moving African continent moves at 3–7 mm y1 NW to WNW and collides with the Eurasian plate (e.g., Serpelloni et al., 2007). Africa has unusually high elevation, in large part owing to the broad plateau uplifts associated with mantle plumes and volcanic construction in the large igneous provinces of the East African Rift System (EARS) and the CVL and the unextended Hoggar and Tibesti volcanic uplifts (Figure 1). Extension effectively increases the surface area of continental landmasses and creates continental platforms. Magmatism during rifting creates new continental crust through dike and sill intrusion and surface construction, but the rates and volumes of crustal accretion are poorly constrained (Lee et al., 2016; Thybo & Artemieva, 2013), in part motivating this comparative study. 2.1. East African Rift System Active faulting and magmatism occur across a large part of the African continent between the Horn of Africa southwest to the Okavango region of Botswana and in a second diffuse arm that continues from southwestern Ethiopia to southeastern Mozambique, including offshore regions of the Indian Ocean, such as the Davie Ridge (Figures 1 and 2). The southern Red Sea, Main Ethiopian, and Eastern, Western, and Southwestern rift systems have developed atop broad topographic plateaux, whereas the Malawi rift and its southeastward continuation transect low elevation regions in Mozambique (Figure 2). The broad plateaus, their corresponding negative Bouguer gravity anomalies, low upper mantle seismic velocities, and the large volume and geochemistry of eruptive volcanic products have been cited as evidence for one or more mantle plumes beneath sections, or all, of the uplifted zones of Africa (e.g., Ebinger & Sleep, 1998; Marty & Yirgu, 1996; Nyblade & Robinson, 1994; Sengör & Burke, 1978). Global and local tomography and geochemical studies reveal low-velocity zones caused by one or a combination of elevated temperatures and the presence of melt in the asthenosphere (e.g., Adams et al., 2012; Debayle et al., 2001; Fishwick & Bastow, 2011; Mulibo & Nyblade, 2013; O’Donnell et al., 2013; Ritsema et al., 1999; Rooney et al., 2011; Simmons et al., 2007; Weeraratne et al., 2003). Although gaps remain in our knowledge of the Western rift and beneath the Indian Ocean, the lowest P and S wave velocity regions underlie the Main Ethiopian Rift (MER) and the Eastern rift zone and the isolated volcanic provinces of

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the Western rift (e.g., Adams et al., 2012; Bastow et al., 2008; Fishwick, 2010; O’Donnell et al., 2013). The Ethiopia-Yemen and East African Plateaux are separated by an ~300 km wide topographic depression that is underlain by crust stretched during Mesozoic rifting, allowing the possibility that the two plateaux are actually one uplifted region extending from southern Africa to the Red Sea: the African superplume province (e.g., Kendall & Lithgow-Bertelloni, 2016; Nyblade & Robinson, 1994; Ritsema et al., 1999). Thermal-mechanical modeling indicates that widespread Mesozoic rift zones spanning the breadth and length of Africa would have been actively subsiding and underlain by thinned lithosphere at the onset of flood magmatism. At ~40 Ma the earliest flood basalts were erupted in southwestern Ethiopia and northern Kenya, suggesting that these thinned and heated regions may have been inherently weaker than surrounding regions (e.g., Hendrie et al., 1994; Morley et al., 1999) or may have ponded anomalously hot mantle material susceptible to decompression melting (e.g., Ebinger & Sleep, 1998). Edge-driven convection may enhance magma production and lithospheric heating at the edges of the deeply rooted Tanzania and Congo cratons where kimberlitic and carbonatitic magmatism has occurred since ~40 Ma (e.g., Ebinger & Sleep, 1998; King & Anderson, 1998; King & Ritsema, 2000). The geochemistry of Eocene-Recent eruptive volcanic products points to a mantle plume origin for the Ethiopia-Yemen flood basalt sequences, but the East African plateau region south of Ethiopia shows more spatial variability (e.g., Chakrabarti et al., 2009; Furman, 2007, Furman et al., 2006; Halldórsson et al., 2014; Pik et al., 2006; Rooney et al., 2011). Although sub-Saharan Africa is widely separated from subducting slabs that carry fluids to the mantle, and the last orogenesis occurred over 500 Myr ago, African volcanoes and active fault systems transfer large volumes of exsolved volatiles, and xenoliths are heavily metasomatized (e.g., Chesley et al., 1999; Frezzotti et al., 2010; Reisberg et al., 2004; Rooney et al., 2017; Trestrail et al., 2017; Vauchez et al., 2005). The lithospheric heating and fluid migration modify crust and mantle density, geothermal gradients, and hydration state, which are the primary controls on the lithospheric strength (e.g., Bürgmann & Dresen, 2008; Hacker et al., 2015; Lowry & Pérez-Gussinyé, 2011). Mineral physics, seismic and magnetotelluric (MT) imaging, and xenoliths provide increasing evidence that hydration state and high partial pressures of CO2 and temperature significantly influence the rheology, density, seismic velocity, and thermodynamics of minerals (e.g., Guerri et al., 2015; Selway et al., 2014; Schmandt & Humphreys, 2010; Vauchez et al., 2005; Wada et al., 2008). Gas and fluid phases also change the frictional properties of fault zones (e.g., Niemeijer & Spiers, 2006). Thus, heat transfer and the migration of magmatic fluids and exsolved gases through the thinning plate beneath rift zones may play a key role in strain localization during early stage rifting (e.g., Buck, 2004; Lee et al., 2016; Maccaferri et al., 2011; Muirhead et al., 2016; Rooney et al., 2017), as outlined in this review paper. 2.2. Cameroon Volcanic Line The CVL formed between the northern edge of the deeply rooted Congo craton and the Benue trough, a highly extended Cretaceous rift system that connected to basins in eastern Sudan and Kenya (e.g., Benkhelil, 1989; Fairhead & Binks, 1991; Poudjom Djomani et al., 1997) (Figure 3). The CVL straddles the continent-ocean boundary (Figure 3). Intriguingly, 42 Ma to Recent volcanism along the CVL displays no clear age progression (e.g., Fitton, 1980; Halliday et al., 1990; Marzoli et al., 2000; Nkouathio et al., 2008), an observation that has prompted numerous workers to explore beyond the classic plate-plume hypothesis for hot spot development to explain the CVL. Alternatives have included decompression melting beneath reactivated shear zones in the lithosphere (e.g., Fairhead, 1988; Fairhead & Binks, 1991; Freeth, 1979; Moreau et al., 1987), small-scale upper mantle convection that may advect mantle lithosphere (e.g., King & Anderson, 1998; King & Ritsema, 2000), and delamination (e.g., De Plaen et al., 2014; Fourel et al., 2013). It has also been suggested that lateral flow of buoyant asthenosphere, beneath continental lithosphere thinned extensively during Mesozoic rifting, may now be contributing to the younger volcanism along the line (Ebinger & Sleep, 1998). In support of this hypothesis, Perez-Gussinyé et al. (2009) constrained lithospheric strength via study of effective elastic plate thickness (Te) across the African continent using coherence analysis of topography and Bouguer anomaly data. Their study revealed corridors of relatively weak lithosphere that continue across the African continent from the Afar region to Cameroon, where Te is also depressed in comparison to the surrounding cratons.

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2.3. Atlas Mountains The Atlas Mountains of Morocco are a 2,000 km long intracontinental compressional belt that strikes ENE from Morocco into Algeria and Tunisia (Figures 1, 3, and 4). The structure was formed by reactivation of Triassic-Jurassic age normal faults that originally formed during the opening of the North Atlantic, followed by the Cenozoic collision of Africa with Eurasia (e.g., Gomez et al., 2000; Pique et al., 2002). Unlike some of the other orogenic belts around the Mediterranean, the origin of the topography of the Atlas is not directly linked to slab rollback (e.g., Dewey et al., 1989; Faccenna et al., 2004; Wortel & Spakman, 2000). The unusually high topography, modest tectonic shortening, thin lithosphere, and localized alkali volcanism may be explained by upwelling of a hot mantle anomaly linked to the Canary island hot spot (e.g., Duggen et al., 2009; Miller et al., 2015; Miller & Becker, 2014).

3. Embryonic Rifting of Cratonic Lithosphere and the Role of Volatiles A long-standing question in plate tectonics involves the initiation of rifting in strong cratonic lithosphere, where the inherent strength of the plate is greater than the far-field and gravitational potential energy forces available (e.g., Bott, 1990; Flesch & Bendick, 2012; Stamps et al., 2014). Where magma is generated, the buoyancy forces of magma add to the tectonic stress, and dike intrusion may accommodate extension at one eighth the force required to overcome friction along a fault (e.g., Bialas et al., 2010; Buck, 2004). Yet only very small melt volumes can be generated beneath thick lithosphere; magma-assisted rifting is unlikely in lithosphere greater than ~ 100 km thick, unless the upper mantle is anomalously hot, carbonated, and hydrated (e.g., Dixon et al., 2008; Turner et al., 1994). The 20 km from the active carbonatitic volcano, Oldoinyo Lengai (Figure 5). The volume of fault zone degassing along faults and from volcanoes may be 11% of the global CO2 budget (Lee et al., 2016). The localized metasomatism, magma intrusion, and magma degassing may have weakened the cratonic lithosphere to enable rift initiation at the edge of the deeply rooted Tanzania craton. Crustal xenoliths from Archaean crust (west) and Pan-African crust (east) are mafic granulites (Jones et al., 1983; Mansur et al., 2014) (Figure 5). Yield strength envelopes for granulitic lower crust indicate that it is too weak to explain the unusual lower crustal seismicity in this rift sector (e.g., Wang et al., 2012; Weinstein et al., 2017). 3.1. Crustal Velocities, Thinning, and Evidence for Intrusions Seismic refraction/wide-angle reflection, arrival time and mantle tomography, MT, and gravity data acquired as part of the KRISP94 (Kenya Rift International Seismic Project) with profile locations shown in Figure 5 provide details of Archaean and Late Proterozoic crust, as well as rift structure. Data from two temporary seismic arrays spanning three rift segments of the Eastern rift, including the active Oldoinyo Lengai volcano, provide

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Figure 5. Compilation of seismic and magnetotelluric (MT) data sets from the Eastern rift system, with respect to crust and mantle xenolith localities. Inset places study area within context of region shown in Figure 2b. Green bold lines show the locations of cross sections of the Natron (A-A0 ) and Magadi (B-B0 ) basins shown in Figure 6. Magadi and Natron are early-stage magmatic rift basins described in the text. Black box encloses the Chyulu Hills off-rift magmatic zone. Black dashed line is the approximate surface contact between Archaean and Pan-African crust. Triangles and inverted triangles are locations of broadband seismic data sets (Albaric et al., 2014; Dugda et al., 2005; Gao et al., 1997; Last et al., 1997; Plasman et al., 2017; Tugume et al., 2012; Velasco et al., 2011; Weinstein et al., 2017), and circles and hexagons are locations of intermediate and long-period MT recordings (Abdelfettah et al., 2016; Hautot et al., 2000; Hautot & Tarits, 2015; Sakkas et al., 2002; Selway et al., 2014; Simpson, 2000). Bold lines are approximate locations of the Kenya Rift International Seismic Project (KRISP) wide angle reflection/refraction and MT profiles as outlined in Prodehl et al. (1994) and Khan et al. (1999).

critical insights into magmatic modification and crustal stretching processes within the faulted basins and beneath the numerous eruptive volcanic centers (Ibs-von Seht et al., 2001; Plasman et al., 2017; Roecker et al., 2017) (Figures 2b and 5). The uppermost crust is dominated by the ~3–7 km thick sedimentary basins that formed in the OligoceneRecent zones of crustal stretching (e.g., Hendrie et al., 1994; Morley et al., 1999; Morrissey & Scholz, 2014). Upper crustal velocities in crystalline basement are 6.0–6.3 km/s. The sometimes-reflective boundary

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between upper and middle crust is marked by a velocity increase to ~6.5 km/s (Birt et al., 1997; Keller et al., 1994). Lower crustal velocities are variable but 6.7–6.9 km/s on average. Last et al. (1997) and Dugda et al. (2005) analyzed receiver functions from widely spaced stations spanning the Archaean-Pan-African suture in the southernmost (Tanzania) and central (Kenya) sectors of the Eastern rift (Figures 2 and 5). Last et al. (1997) find no change in crustal thickness crossing the Archaean-Proterozoic boundary through the central Manyara rift. Dugda et al. (2005) find slightly thicker crust in the Pan-African belt east of the rift (39–42 km) than west of the rift beneath the Tanzania craton (37–38 km), and normal Poisson’s ratios of 0.24–0.27, indicating little or no crustal thinning or magmatic modification outside the rift zones. Geophysical data from the Magadi-Natron-Manyara sector of the Eastern rift reveal crustal thinning and magma intrusion across the rift zone and low-velocity, high VP/VS bodies beneath some volcanoes. Moho depth determined from receiver functions and the KRISP94 profile varies between 28 and 41 km beneath the rift valley and rift shoulders, respectively (Figures 2b, 6a, and 6b). The thickest crust (41 km) underlies the Crater Highlands, a wedge of uplifted basement, and 125 km outside the rift zone (Green et al., 1991; Ritsema et al., 1999), consistent with P-T conditions of eruptive volcanic products (Macdonald, 1994; Mechie, Fuchs, et al., 1994). Along the length of the rift, velocities in the mantle just below the Moho from Pn analyses are 7.5–7.7 km/s, significantly lower than values beneath the flanks and stable interior: 8.0–8.1 km/s (Maguire et al., 1994; Mechie, Keller, et al., 1994). The magnitude of the velocity decrease is 0.3 km/s larger than predicted from higher geothermal gradients, suggesting that partial melt within the upper mantle also contributes to the observed velocity reduction beneath the rift (Mechie, Fuchs, et al., 1994). If we assume an initial crustal thickness of 40 km, the ~28 km thick crust beneath the Natron and Magadi basins crustal thickness estimated from wide-angle refraction/reflection studies and receiver function studies indicates ~30% extension achieved in 6.8 km/s have been interpreted as evidence for gabbroic crustal intrusions (e.g., Keranen et al., 2004; Mackenzie et al., 2005; Maguire et al., 2006). In the upper crust, 3-D controlled source tomography shows significant variation in VP, with discrete high VP (>6.4 km/s) zones beneath the magmatic segments (Figures 8 and 9). The >7 km/s velocities and high densities inferred from forward and inverse models of gravity data indicate that the magmatic segments are underlain by gabbroic material (e.g., Cornwell et al., 2006; Mahatsente et al., 1999; Tiberi et al., 2005). The observed pattern of segmented high-velocity zones was reproduced by gabbroic bodies that rise to 10 km below the surface, indicating that 30% of the crust beneath the magmatic segments is new igneous material (Keranen et al., 2004). The observations corroborate the surface geology evidence that axial magma intrusion, likely in the form of gabbroic cumulates from crustal magma chambers, fed sheeted dykes. The dikes, as well as faults above the dikes, accommodate most of the deformation across the rift, based on time-averaged and active deformation patterns (Casey et al., 2006; Keir et al., 2015). New GPS data confirm that active deformation occurs across the Ethiopian Plateau (Birhanu et al., 2016), which hosts aligned chains of Quaternary eruptive centers, highconductivity lower crust, and lower crustal seismicity (Keir et al., 2011a, 2011b) (e.g., Figure 10). 4.2. Crustal Magma Plumbing Systems The interpretation of gravity and electrical conductivity anomalies has been used to construct an ever more detailed picture of the crustal magma plumbing system accommodating extension during breakup and

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Figure 9. Rift cross sections comparing several geophysical properties of the crust across the MER. The position of C-C0 is labeled on Figure 7a. (a) The 2-D resistivity structure of the crust along a portion of A-A0 determined using MT methods (Whaler & Hautot, 2006). The shallow low-velocity parts of the model are interpreted as sediments and volcanic and volcaniclastic rocks filling in the subsiding basins of the MER, extending between ~ 150 and 250 km of C-C0 (Keranen et al., 2004; Mackenzie et al., 2005). (b) Profile C-C0 is the P wave velocity model of the crust determined using controlled source reflection/refraction (Mackenzie et al., 2005) with earthquake hypocenters (gray dark circles) recorded during October 2001 to January 2003 and located within 20 km either side (dashed lines) of the profile projected onto the section. Labels are u.c., upper crust; l.c., lower crust; HVLC, high-velocity lower crust; M, Moho.

sourcing active off-axis volcanoes (Figure 9). In the MER, Cornwell et al. (2006) conducted a cross-rift gravity survey and Whaler and Hautot (2006) a cross-rift MT survey, coincident with wide-angle Line 1 (Figure 9). Major features of the gravity study are an axial low-density upper mantle or high-density lower crustal zone which is modeled as a ~50 km wide body with a density of 3,190 kg/m3, supporting interpretations of a mafic underplate layer beneath the northwestern rift flank (Cornwell et al., 2006). Two high-density (3,000 kg/m3) upper crustal bodies underlie the MER: a 20 km wide axial body and a 12 km wide off-axis body, both of which are likely gabbroic in composition. MT data provide some of the best evidence for the emplacement and migration of melt during breakup. Major features of the MT survey are a conductive body at 20–25 km depth beneath the rift axis interpreted as a middle-to-lower crustal magma reservoir (Figure 9). A second highly conductive region at 25–35 km depth is beneath the rift flank near the SDFZ and YTVL (Figures 6 and 8). The lower crustal high conductivity zone coincides with seismicity, leading to its interpretation as a pressurized magma reservoir beneath the rift flank volcanoes (Keir et al., 2009) (Figure 9). Ambient noise tomography reveals low VS (2, implying that melts are also present (Dugda et al., 2005; Hammond et al., 2011; Stuart et al., 2006). However, Hammond (2014) shows that such high values are an effect of seismic anisotropy, suggesting that if melt is the cause, then it must be preferentially aligned. Inverting the azimuthal variations in

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VP/VS suggests that melt in the lower crust must be stored in interconnected sills, consistent with reflectivity studies in rift zones (e.g., Thybo et al., 2000). Estimates of azimuthal anisotropy for the whole crust have the same orientation as that in the upper crust but with much stronger anisotropy. Further support for aligned crustal melt pockets comes from measurements of geoelectrical strike orientations determined from the MT profiles outlined in section 5.2 (Johnson, 2012). The presence of electrical anisotropy implies that the aligned melt is interconnected. For periods synonymous with the mostly low-conductivity upper crust (1–10 s), the electrical anisotropy strikes are similar to those from upper crustal earthquake shear wave splitting, and the magnitude of electrical anisotropy (measured as the difference in MT phase for currents flowing along and across geoelectrical strike; see, e.g., Hamilton et al., 2006; Padilha et al., Figure 12. (a) Variation in Quaternary-Recent basaltic volcanism exposed 2006) is relatively low. For periods synonymous with the conductive lower every 0.1° latitude along strike in the Red Sea rift in Afar, with crustal velocity structure largely based on the EAGLE along-axis seismic profile (Maguire crust (100–1,000 s) the strike of the electrical anisotropy matches well with et al., 2006) and location shown in Figure 7a. (b) Elevation and (c) crustal fast directions from receiver functions and teleseismic shear wave splitthickness. Seismic velocities are in km/s. Note the abrupt thinning of the ting, and the amount of electrical anisotropy is higher (Figure 11). These o crust in northern Afar (>13 N), which coincides with subsidence of the rift similarities support the idea that seismic and electrical anisotropy are valley below sea level and a marked pulse in Quaternary-Recent volcanism caused by the preferential alignment of melt, primarily in the lower crust. (after Bastow & Keir, 2011). SDR = seaward dipping reflector; δVP indicates regional velocity reduction relative to global norm. These patterns suggest that melt can flow within these interconnected networks in the lower crust and feed the relatively localized zones of deformation in the upper crust (Desissa et al., 2013; Hammond, 2014). Note that there are areas with little or no electrical anisotropy, but high conductivities indicating abundant melt, in the upper crust (e.g., close to Dabbahu volcano; Figure 11), suggesting no particular geometry in its arrangement. Here the penetration depth is severely limited, and the 100–1,000 s periods are probably not penetrating the deeper crust. 5.4. Crustal Thinning and Subsidence at Plate Rupture Wide-angle seismic experiments show that the crust throughout Ethiopia has a consistent layering with lower crustal VP = 6.7–7.0 km/s, an upper crustal VP = 6.0–6.3 km/s, and cover rocks of lava flows and sediments with VP = 2.2–4.5 km/s (Makris & Ginzburg, 1987; Prodehl & Mechie, 1991a, 1991b) (Figures 8, 10, and 12). Excluding the regularly spaced volcanoes marking the centers and tips of magmatic segments, the cover rocks are generally thickest where the crust is thinnest suggesting a strong link between crustal thinning, rift valley subsidence, and resultant accumulation of basin infill (Figure 12). The Danakil basin, which hosts the Erta’Ale magmatic segment, has the thinnest crust in Afar at 15 km thick and is also characterized by the thickest sedimentary basins: ~5 km as opposed to 2–3 km elsewhere (Berckhemer et al., 1975; Makris & Ginzburg, 1987) (Figure 12). The change in crustal thickness along rift from south to north is primarily accounted for by markedly thinned lower crust, although the upper crust is also appreciably thinner (Makris & Ginzburg, 1987). Existing controlled source data do not traverse the rough relief of the currently active magmatic segments where intrusion volumes may be largest. Spatial variations in crustal thickness in Ethiopia correlate with variations in effective elastic thickness, with the strongest plate (Te ≈ 60 km) beneath the plateaus and the weakest plate (Te ≈ 6 km) beneath the Danakil depression (Perez-Gussinyé et al., 2009). Both crustal thickness and effective elastic thickness, a measure of strength over time periods much longer than the earthquake cycle, also correlate to variations in seismogenic layer thickness estimated from local seismicity and teleseisms; seismogenic layer thickness is 25–30 km beneath the Ethiopian Plateau (Keir et al., 2009) but decreases to ~5 km beneath the Danakil depression (Craig et al., 2011; Nobile et al., 2012). There is debate regarding whether the regions of thinnest crust (~15–20 km) in Afar, such as the Danakil depression, are fully oceanic in nature. In the Danakil depression, the basin floor is at ~120 m below sea level and the Holocene stratigraphy is dominated by intercalated basalts and evaporites (Atnafu et al., 2015). The basin stratigraphy and environment of deposition are similar to that interpreted to form the seaward dipping reflector sequences (at fully rifted margins), and therefore, the Danakil depression could be a modern analog for their formation (Bastow & Keir, 2011; Buck, 2017; Corti et al., 2015). There is at least 50% more melt available

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beneath the northern MHR than in the TG segment in the south and likely up to an order of magnitude more. This change is consistent with regional northward increases in the number of Holocene volcanoes as well as the volume of erupted material (Barberi & Varet, 1977), and it is likely a result of a northward increase in thinning of the plate and resultant decompression melting of asthenosphere (Bastow & Keir, 2011). The most highly extended crust beneath the Danakil depression is 5–10 km thicker than normal oceanic crust. There are two end-member hypotheses to explain the structure and composition of the Danakil crust. In one model, heavily intruded and extended crust deforms by brittle failure at the surface and by magma intrusion and ductile flow in the lower crust (Bastow & Keir, 2011). Unlike thin crust at oceanic rifts, the intrusive material may stall and fractionate to produce peralkaline and rhyolitic lavas, and felsic intrusive material, contributing to the unusual structure. The second model relates to high magma production rates in earlystage seafloor spreading caused by the steep lithosphere-asthenosphere boundary on either side of the rift (e.g., Boutilier & Keen, 1999; Korenaga et al., 2000; Ligi et al., 2011; White & McKenzie, 1989). The high extrusion rates at the weak spreading ridge cause bending above the ridge and may produce seaward dipping reflectors. Our syntheses suggest that both processes are important. 5.5. Summary: Defining Transitional Crust and Plate Rupture Along Magmatic Margins Passive and controlled source seismic imaging shows that the crust in Afar varies from 15 to 25 km thick, and the crust is heavily modified by solidified mafic intrusion and partial melt currently residing in magma reservoirs through the crust. MT combined with modeling seismic and electrical anisotropy supports the interpretation that the lower crust is an important region of melt ponding. The melt in the lower crust resides in horizontal sills near segment centers, and it is then transported laterally and vertically in dikes that accommodate most of the plate boundary deformation. The central magma chamber that maintains the along-axis segmentation over multiple episodes and the extension via dike intrusion and faulting above the dikes produce columns of new crust with corresponding magnetic stripes. The unusually thick crust (~15 km) matches the thick crust beneath magmatic rifted margins with seaward dipping reflector sequences, pointing to a common mode of emplacement. In contrast, the observed seismic velocity and structure of the crust and the low-silica composition of basalts suggests that seafloor spreading processes are still evolving. Seismic imaging, gravity, and MT all suggest that the whole crust beneath the rift flanks and the rift axis is heavily modified by magma intrusion. Seismic and electrical anisotropy support the view that magma accumulates in the lower crust as networks of interconnected stacked sills, whereas magma in the upper crust may be both in sill networks and subvertical dikes. Petrological constraints suggest that fractionation starts in the lower crust beneath flank volcanoes during the early stages of rifting, whereas most fractionation occurs in the upper crust beneath axial volcanoes as plate rupture initiates (Rooney et al., 2011, 2014).

6. The Cameroon Volcanic Line Cameroon’s crustal structure has been investigated using both passive and active seismic sources (e.g., Dorbath et al., 1986; Gallacher & Bastow, 2012; Meyers et al., 1998; Plomerova et al., 1993; Stuart et al., 1985; Tokam et al., 2010) and potential field studies (e.g., Fairhead & Okereke, 1987; Tadjou et al., 2009; Toteu et al., 2004) (Figures 3b and 13). Thinned crust (∼25 km) resulting from extension is imaged beneath the Garoua rift (Figure 12), while the transition from the CVL to the Congo craton is heralded by a transition in Moho depth from ∼36 km to 43–48 km depth (Gallacher & Bastow, 2012; Tokam et al., 2010) (Figure 13). Crustal shear wave velocities also increase into the craton from the CVL, from ∼3.7 km/s to ∼3.9 km/s (Tokam et al., 2010). Bulk crustal VP/VS ratios along the CVL are comparable to those of cratons worldwide (∼1.74), an observation that Gallacher and Bastow (2012) cited as evidence for the lack of melt fractionation and intrusion during the ∼30 Ma development of the CVL. These seismological results thus corroborate petrological studies that attribute low-volume, high-pressure magmas to melting of subcontinental lithospheric mantle that has experienced only small amounts of crustal fractionation (e.g., Fitton, 1980; Halliday et al., 1990; Marzoli et al., 2000; Suh et al., 2003; Yokoyama et al., 2007). Such low-volume, high-pressure magmas are expected to form within the subcontinental lithospheric mantle and exhibit relatively little fractionation within the crust (e.g., Suh et al., 2003). Gravity studies in Cameroon (e.g., Fairhead & Okereke, 1987 ; Poudjom Djomani et al., 1992, 1997; Tadjou et al., 2009) highlight various major tectonic features in the region, including the Benue Trough (+ve

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Figure 13. (a) Location map of the CBSE seismograph stations (triangles) superimposed on regional topography. Numbers are station codes. The Ntem Complex boundary and the Yaoundé domain boundary are after Toteu et al. (2004). Stars are selected CVL volcanoes. C.A.R.: Central African Republic; CASZ: Central African Shear Zone; E.G.: Equatorial Guinea. The red areas are regions of Cenozoic volcanism along the CVL taken from Tokam et al. (2010). (b) Variations in crustal thickness. (c) Variations in VP/VS ratio across Cameroon from receiver function analysis. (d) The black thick lines A-A0 , B-B0 , and C-C0 show the orientation of transects. After Gallacher and Bastow (2012).

Bouguer anomaly), uplift of the Adamawa Plateau (ve anomaly), and the Congo craton (ve Bouguer anomaly). Evidence for continental collision was found at the northern margin of the Congo craton, as well as a positive anomaly associated with the Central African Shear Zone, which bisects the Adamawa uplift (Figure 13, e.g., Poudjom Djomani et al., 1997; Tadjou et al., 2009). Recent work by Milelli et al. (2012) and Fourel et al. (2013) has used scaled laboratory models and analytical solutions to investigate the effects of large lateral variations in lithospheric thickness on lithospheric stability. Their work demonstrates that lithospheric instabilities can develop over long timescales with small rates of

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Figure 14. S receiver function profile across Atlas Mountains from northwest to southeast as indicated in Figure 3a. Top panel indicates the elevation and primary tectonic features (HMP-High Moulouya platform; FMA-folded Middle Atlas; TMA-Tabular Middle Atlas). S receiver function stacks at 40 km spacing are shown for all stations within 20 km of profile and plotted evenly spaced for clarity. Blue and red dashes indicate Moho and lithosphere-asthenosphere boundary, respectively.

upwelling and decompression melting as is observed in Cameroon and consistent with recent hybrid models for magmatism along the CVL (e.g., De Plaen et al., 2014; Gallacher & Bastow, 2012; Reusch et al., 2010).

7. The Atlas Orogen In contrast to the Mediterranean arcs (Alps, Apennine, Carpathian, Hellenic, and Betic-Rif), the Atlas is subparallel to and south of the convergent margin and on the African plate (Figures 1 and 3a). The mountain belt has high topography (>4,100 m), yet there is no apparent deep crustal root to isostatically support the high elevations of the Atlas (e.g., Ayarza et al., 2005, 2014; Jessell et al., 2016; Miller & Becker, 2014; Missenard et al., 2006; Sandvol et al., 1998; Zeyen et al., 2005) (Figure 3a). Furthermore, only a modest amount of tectonic shortening has been estimated and has been suggested to be achieved through thick-skinned thrusting and folding (e.g., Gomez et al., 2000; Teixell et al., 2003). Many recent efforts have focused on understanding the crustal- and lithospheric-scale structure of the Atlas to understand the orogenesis. Structural seismological imaging suggests that the lithosphere beneath the Atlas is particularly thin or perhaps the uppermost mantle is abnormally warm with low seismic velocities found at depths of ~65–160 km (e.g., Bezada et al., 2014; Fullea et al., 2007; Miller et al., 2015; Palomeras et al., 2014; Sun et al., 2014). Figure 14 shows S receiver function estimates of the Moho and lithosphereasthenosphere depths along a profile across the Atlas from Miller and Becker (2014). Geophysical modeling indicates that the lithosphere is thin (~65 km) compared with the topographically high, thick lithosphere of the Saharan platform (≥150 km) and the Morocco Atlantic margin along a corridor that is aligned with the highest topography in the Atlas (Fullea et al., 2007; Miller & Becker, 2014; Miller et al., 2015; Missenard et al., 2006; Missenard & Cadoux, 2012; Teixell et al., 2003, 2005) (see Figure 3a). The combination of thinned lithosphere and mantle upwelling has been invoked to explain the unusual high topography, low seismic velocities in the uppermost mantle, and lack of a significant crustal root (Arboleya et al., 2004; Frizon de Lamotte et al., 2009; Fullea et al., 2007; Miller & Becker, 2014; Missenard et al., 2006; Teixell et al., 2005; Zlotnik et al., 2014). The source of the upwelling has been suggested to be part of the Canary plume from geochemical analyses of Quaternary alkali basalts in the Atlas and their spatial location above the low velocity material beneath the mountain range (Anguita & Hernán, 2000; Duggen et al., 2009). Shear wave splitting analyses (Díaz & Gallart, 2014; Miller et al., 2013) also indicate a strong change in azimuthal anisotropy strength at the northern edge of the Middle Atlas, as well as an alignment of fast polarization orientations parallel to the strike of the Atlas Mountains, suggesting shearing in a mantle channel guided by lithospheric topography (Miller et al., 2013; Miller & Becker, 2014) and deflected by the high-velocity slab

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beneath the Alboran Sea (Alpert et al., 2013; Díaz & Gallart, 2014). Miller et al. (2015) interpret the localized volcanism and low velocity anomalies in the sub-lithospheric mantle to be the result of the Canary Island plume flowing sub-horizontally beneath the Atlas. This seismologically based interpretation is supported by numerical experiments of mantle flow that incorporate the effects of the stiff, deep West African cratonic keel and nearly vertical, narrow slab beneath the Alboran (Alpert et al., 2013), numerical experiments of strongly tilted plumes and resulting instabilities (Mériaux et al., 2011), and by scaled analogue models of the Alboran slab and Canary plume interaction (Mériaux et al., 2015). Other hypotheses explaining the Cenozoic uplift of the Atlas and the links to recent volcanism and to upper-mantle structure include: lithospheric delamination (Bezada et al., 2014; Duggen et al., 2009; Levander et al., 2014) and edge-driven convection (Kaislaniemi & van Hunen, 2014; Missenard & Cadoux, 2012). The two edge-driven convection studies (Kaislaniemi & van Hunen, 2014; Missenard & Cadoux, 2012) are based on numerical experiments, which are designed to evaluate the dynamics of the thinned lithosphere and mantle melting processes in the Moroccan Atlas. They also find good agreement with observations of lithospheric thickness and volcanism, but neither incorporate the presence of the subducted slab beneath the Alboran Sea. Despite the lack of consensus on the origin of the Atlas topography, various recent studies using a range of methods, from the SIMA active source seismic experiment (Ayarza et al., 2014), broadband waveform analysis (Sun et al., 2014), MT (Anahnah et al., 2011; Ledo et al., 2011), body wave (e.g., Bezada et al., 2014), and surface wave (Palomeras et al., 2014) tomography, for example, have suggested the presence of anomalous lowvelocity material beneath the lithosphere and in the crust. The temperature of this material is also inferred to be warm, due to the presence of partial melt (Anahnah et al., 2011; Ayarza et al., 2014; Ledo et al., 2011; Sun et al., 2014).

8. Summary Our synthesis of geophysical studies that aim to image the crust beneath zones of active rifting and orogenesis in Africa shows that magma and volatiles are migrating from the asthenosphere through the plate, leading to changes in rheology and making large contributions to H2O and CO2 fluxes globally. The addition of magmatic fluids permanently alters the composition of continental crust accreted during the Proterozoic, and it increases the continental crustal volume. The magma intrusion accommodates plate boundary deformation and masks crustal stretching. In young, weakly extended sectors of the East African rift, volatiles released from upwelling asthenosphere, heated mantle lithosphere, and magma intrusions alter crustal composition and rheology. Even in rifts that formed within the last 7 Myr, ~20% of the crust is new magmatic material intruded as sills in the lower crust (underplate) and as more localized zones of dike intrusion in the middle and upper crust as imaged with seismic, MT, and gravity methods. In the mature Main Ethiopian Rift (MER), the surface area of the rift has doubled in the past ~ 10 Myr, and intrusive volumes are 30%, increasing into the central Afar rift zone where dike intrusion accommodates much of the active extension, and symmetric magnetic anomalies mimic seafloor spreading. Intrusive to extrusive volumes (surface flows plus basin fill) are roughly equivalent but increase with increasing age of rifting and amount of extension and proximity to prerift flood magmatism. In the older, more evolved Horn of Africa, intense episodes of magma intrusion fed from crustal magma chambers beneath segment centers create new columns of mafic crust up to 8 m wide in a process akin to that along slow-spreading ridges. The magma intrusion and faulting above the dikes accommodate centuries of plate boundary opening and suggest that plate rupture is achieved through catastrophic magma intrusion events in 15–20 km thick transitional crust. Regional analyses suggest that progressively smaller screens of continental crust are heated and weakened by repeated episodes of magma intrusion immediately prior to seafloor spreading. On the western side of the African continent, magmatism with very minor extension over the past 30 Myr has formed the Cameroon Volcanic Line (CVL)—a linear chain of basaltic volcanoes that formed on the broad Adamawa Plateau that lacks the age progression predicted by the traditional hot spot track hypothesis. Intrusive volumes are much smaller than in the Eastern, MER, and Afar rifts, and magmas largely fractionate at subcrustal depths. The comparison with East African crustal structure suggests that extensional forces are much smaller in the CVL, inhibiting the creation and rise of magma into the crust and preventing full-scale rifting. In the Atlas orogen of northwestern Africa, recent magmatism is also primarily along a linear trend from the Canary Islands to the Alboran Sea and linked to the Canary Island hot spot. In contrast to the

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magma-rich Horn of Africa, however, magmatism in Cameroon and Morocco has had relatively little impact on crustal structure, but the mantle flow and lithospheric heating cause dynamic uplift supporting the high Atlas mountain range.

Acknowledgments We thank two anonymous referees and Editors Jeff Gu and John Geissman for feedback that greatly improved text and figures. C. E. was supported by National Science Foundation grants EAR-1113355 and EAR-1109302. D. K. is supported by Natural Environmental Research Council grant NE/L013932/1. C. E. and D. K. also were supported by a grant from Beach Petroleum and Tanzania Development Corporation. C. T. and S. H. were supported by ANR grant 12-JS06-000. C. E., C. T., and D. K. conducted studies with approval of the Commission for Science and Technology (Tanzania) and the National Council for Science and Technology (Kenya). K. A. W. was supported by NERC grants NER/B/S/2001/00863 and NE/E007147/1 and is currently supported by NERC grant NE/L013932/1. Her MT fieldwork used equipment loaned by the NERC Geophysical Equipment Facility and from the University of Brest and the Geophysical Instrument Pool Potsdam and was supported logistically by Addis Ababa University, the Geological Survey of Ethiopia, and the Petroleum Operations Department of the Ethiopian Ministry of Mines. M. S. M. is supported by NSF Continental Dynamics Program EAR0809023 and NSF CAREER award EAR1054638. The facilities of SEIS-UK are supported by the Natural Environment Research Council under agreement R8/H10/64. The facilities of IRIS Data Services, and specifically the IRIS Data Management Center, were used for access to waveforms, related metadata, and/or derived products used in this study. Broadband seismic data collected in Morocco utilized PASSCAL instrumentation and were supported by IRIS. IRIS Data Services are funded through the Seismological Facilities for the Advancement of Geoscience and EarthScope (SAGE) Proposal of the National Science Foundation under Cooperative Agreement EAR-1261681. CRAFTI data (XJ) are available at the IRISDMC in 2017 (http://ds.iris.edu/ds/ nodes/dmc/).

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Our work highlights consistent patterns and new insights into continental lithospheric deformation processes. Based on our results and complementary crust and mantle xenolith studies, future studies would benefit from constraints on physical properties across a broader range of crustal and magmatic rock compositions, and in particular, relatively low silica, and gas-rich melts. Our work focuses on areas with sometimes dense data coverage, but large sectors of the African continent remain virtually unexplored in terms of crust and upper mantle structure through seismic, MT, heat flow, and detailed structural analyses. Gap-filling studies will tighten the context for studies of active deformation, and they promise new insights into continental lithospheric behavior, given the relative stability of the African continent. Critical areas include the largely amagmatic but remote Western and Southwestern rift zones, the Turkana gap between the Ethiopian and East African Plateaux, the seismically active Indian Ocean margin, and the boundaries of thick Archaean cratons, which have been the locus of plate deformation through multiple Wilson Cycles.

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